Sustained and intensified lacustrine methane cycling during Early Permian climate warming

Lakes are a major emitter of the atmospheric greenhouse gas methane (CH4); however, their roles in past climate warming episodes remain poorly understood owing to a scarcity of geological records. Here we report the occurrence of sustained and intensified microbial CH4 cycling in paleo-Lake Junggar in northwestern China, one of the largest known Phanerozoic lakes, during Early Permian climate warming. High-precision U-Pb geochronology refines the age of the upper Lucaogou Formation to the Artinskian, which marks a major glacial-to-postglacial climate transition. The 13C-enriched authigenic dolomites indicate active methanogenesis in the anoxic lake sediments, and 13C-depleted hopanes suggest vigorous methanotrophy in the water column. The intensification of CH4 cycling coincided with increasing global temperature, as evidenced from elevated continental chemical weathering. Our results suggest that the lacustrine CH4 emissions acted as a positive feedback to global warming and contributed to the demise of the Late Paleozoic Ice Age.

interbedded within the black organic-rich shales ( Fig. 2a; Supplementary Fig. 1g-i). They are microcrystalline (typically < 10 µm) and contain dolomite microspheres ( Supplementary Fig.   2e,f), which were interpreted as methanogen microfossils 2 . Evidence for evaporite minerals and subaerial exposure is absent. Thus, it is concluded that the upper member is the basinal facies deposited in a semi-deep to deep lake environment.

Zircon U-Pb geochronology
In the field, the volcanic ash (sample VA-1) is yellowish-white and occurs as a thin The tuffaceous siltstone (sample TS-1; Supplementary Fig. 1d) 3a). At the top of the succession, the isotopic values return to higher values of approximately -27‰, comparable to those in the pre-CIE interval. The δ 13 CTLE and δ 13 CAsph values of the total lipid extract (TLE) and asphaltene (Asph) vary from -28.9 to -33.4‰ and -27.8 to -32.4‰, respectively, slightly lower than the corresponding δ 13 Corg records. They also exhibit a largely similar negative CIE to the δ 13 Corg records throughout the succession  Supplementary Fig. 5b). There is only a weak correlation between the δ 13 Corg and TOC values ( Supplementary Fig. 5c).

Dolomite C and O isotopes
The dolomite δ 13 C values in the lower member are characterized by little variations, from +5.3 to +8.3‰ ( Supplementary Fig. 6), and are slightly positive than those recorded in the Early Permian marine carbonates from the Yangtze Platform, South China (ca. 0 to +6‰) 6

Molecular compositions and compound-specific C isotopes
The samples commonly contain a suite of n-alkanes from C17 to C33 without an oddeven C number preference. The C19-C23 n-alkanes are the predominate compounds, with a maximum at C21 in most samples ( Supplementary Fig. 7a). The major isoprenoids are pristane (Pr) and phytane (Ph), with more abundant Ph (Pr/Ph ratio typically ranging between 6 0.4 and 0.9). The samples also include small amounts of steranes. Both C30 17α,21β-hopane and C29 17α,21β-norhopane appear to be the dominant hopanes, together with minor amounts of 17β,21α(H) isomers. The C31-C34 homohopane abundances decrease with an increasing carbon number. Tm (17α-22,29,30-trisnorhopane) and Ts (18α- 22,29,30-trisnorneohopane) are also present, with Ts/(Ts + Tm) ratios ranging from 0.24 to 0. 45 and -54‰ for the C30 and C29 17α,21β-hopanes, respectively (Fig. 3c). The 17β,21α(H) isomers, i.e., C30 and C29 17β,21α-hopanes also have relatively 13 (Fig. 3d). The CIW values tend to be slightly higher than the CIA values, with a similar temporal trend to CIA. The Ti/Al ratios are uniform throughout the stratigraphic section, with an average value of 0.05 ± 0.003 (1 SD; Supplementary Fig. 9c), indicating that the shifts in the CIA and CIW cannot be attributed to changes in sediment provenance. Thus, the profiles of the weathering indices (CIA and CIW) provide reliable records of the variation in paleoweathering in the studied area. 8 According to the CIA-derived temperature estimates 9 (Fig. 3d). However, even if the absolute calculated temperatures are not strictly constrained, we suggest that the shift in climate warming recorded by our enhanced chemical weathering indices (CIA and CIW) is reliable and can be compared with the CIA trends in many contemporaneous successions (as shown in Fig. 2 in ref. 9 ), such as that from the Karoo Basin in South Africa 10 . In addition, the decrease in the δ 18 O values, as recorded by both lowand high-latitudinal fossil shells composed of low-Mg calcite, suggest an increase in the seawater temperature during the Artinskian 11,12 . Thus, the consistent terrestrial and marine temperature estimates indicate that climate warming was prevalent at that time (i.e., the Artinskian Warming Event 13 ).  14,15 . In addition, it has been proposed that the two sources from terrestrial versus lacustrine organic matter show distinctive δ 13 Corg, HI, and C/N signatures 15,16 ; crucially, however, there is no apparent relationship between δ 13 Corg and the HI and C/N ratios in the Lucaogou shales, despite a poor correlation with TOC ( Supplementary Fig. 5a-c). Therefore, it is unlikely that proportional changes in the organic matter from terrestrial versus lacustrine environments also resulted in the observed CIE. For 9 further verification, we report compound-specific C isotope analyses of n-alkanes with different chain lengths (Supplementary Dataset 4). These compounds were derived, in a large part, from the cracking of parent molecules in kerogen. The results show that all the δ 13 Cnalkane records of short-chain n-C19, mid-chain n-C21, and long-chain n-C27 alkanes display a negative CIE with a magnitude of ~4‰, which is similar to the bulk δ 13 Corg record ( Supplementary Fig. 4a). Under our high-precision CA-ID-TIMS age constraint, this negative CIE is consistent with that recorded in coeval marine brachiopod shells from the U.S.  Fig. 4b), and this may have been superimposed on the regional signal linked to isotopic refractionation by microbial chemosynthetic processes 20 .
Crucially, in some lake settings (e.g., Lake Rotsee 21 and Green River Formation 24 ), hopanoid δ 13 C values are strongly 13 C-depleted (lower than -40‰) and up to 10‰ more negative relative to the co-occurring bulk organic matter or n-alkanes (ca. -30‰). In such situations, it is clear that active methanotrophy occurred in the water column and the hopanoids were largely derived from aerobic methanotrophic bacteria. In contrast, hopanoid δ 13 C values higher than -40‰ suggest a dominant contribution from heterotrophic bacteria (e.g., Ace Lake from Antarctica 28 and lacustrine sequences from the Jianghan Basin 25 ), which can also produce hopanoids and are related to relatively positive C isotopic signatures 34 . The data compilation suggests that -40‰ can be considered a baseline for evaluating the substantial contributions of methanotrophs ( Supplementary Fig. 8). Based on a C isotopic mass-balance calculation (see Methods) 21